Salby M.L. Fundamentals of Atmospheric Physics

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540

17 The Middle Atmosphere

iii

1011

I I I I I III I I I I I I Illi I I I I I III

- ~p= 90

(Eo3= 250 DU) -

.................... (p=59 ~

[ oo"

-- \

\ o~ ..... .o _

\ ,o~

- \

\ ,

I I 'F I lill I I I i ill I I I I

5 1012 5 1013

OZONE NUMBER DENSITY (cm 3)

(•03= 450 DU)

Chapman

Chemistry

I III

5 10

Figure

17.2 Vertical profile of ozone number density under photochemical equilibrium, as

calculated from Chapman chemistry (solid line), and observed in tropical (dashed line) and ex-

tratropical (dotted line) regions. Chapman chemistry was calculated with rate coefficients from

Brasseur and Solomon (1986) and Nicolet (1980). It yields realistic vertical structure, but a col-

umn abundance of Eo3 ~ 1000 DU. Observed data are from Hering and Borden (1965) and

Krueger (1973).

the ozone column is concentrated, the photochemical lifetime is long enough

for 03 to be transported by the circulation and therefore to pass between

widely differing photochemical environments. Even at higher altitudes, other

species participating in the complex photochemistry of ozone are influenced

importantly by transport.

17.2 Involvement of Other Species

Photodissociation by UV radiation produces a number of free radicals from

less reactive reservoir species of tropospheric origin. Photochemically active,

these free radicals can then go on to destroy ozone in catalytic cycles that

leave the free radical unchanged.

17.2 Involvement of

Other Species

541

17.2.1 Nitrous Oxide

The free radical nitric oxide, NO, represents an important link to human

activities. Nitric oxide is a by-product of inefficient combustion (e.g., in aircraft

exhaust). In the stratosphere, the principal source of NO is dissociation of

nitrous oxide

N20

through reaction with atomic oxygen:

N20 + O --+

2NO. (17.8)

Nitrous oxide is produced in the troposphere by natural as well as anthro-

pogenic means (Chapter 1). Away from the surface,

N20

is long-lived, having

a photochemical lifetime of order 100 years (Fig. 17.3). Further, nitrous oxide

is not water soluble, so it is immune to normal scavenging mechanisms asso-

ciated with precipitation (Sec. 9.5.2). These properties allow

N20 to

become

well-mixed in the troposphere and make it useful as a tracer of air motion in

the stratosphere.

After entering the stratosphere,

N20 is

photodissociated according to the

reaction

N20 +

--~ N 2 + O,

(17.9)

which constitutes the primary destruction mechanism for nitrous oxide and

is responsible for the decrease of its mixing ratio rN2 o above the tropopause

(Fig. 1.20). The mixing ratio of N20 is largest above the tropical tropopause

(Fig. 17.4), which, in view of its long lifetime, illustrates how air enters the

stratosphere. A dome of high values reflects upwelling motion in the tropical

stratosphere that carries nitrous oxide across the tropical tropopause from its

reservoir below.

Figure 17.3

(1986).

80 ,

E 6O

_ 4O

100

1 min 1 hour 1 day 1month 1year

~' /"

''"11"

'"' '\~"'"i'"

I I I I I I i

Ill Ill ill 111 III ill

101 102 103 104 105 106 107 108 109

LIFETIME (s)

Photochemical lifetimes of nitrogen species. Source: Brasseur and Solomon

542

17 The Middle Atmosphere

rN2 o (ppbv)

100 70

"T" 150

-30 -15 0 15 30 45 60 70

LATITUDE (Degs)

Figure 17.4 Zonal-mean mixing ratio of nitrous oxide N20 on January 9, 1992, as observed

by the ISAMS instrument on board the Upper Atmosphere Research Satellite (UARS). Adapted

from Ruth

et al.

(1994).

The free radical NO, once produced via (17.8), can efficiently destroy ozone

through the catalytic cycle

NO +

03 ~

NO 2 -k-

02,

(17.10.1)

NO 2 -k- O --+

NO +

02,

(17.10.2)

which has the net effect

03 -1- 0 ~ 202

(net). (17.10.3)

This closed cycle leaves NO + NO2 unchanged, so one molecule of nitric oxide

can destroy many molecules of ozone.

Reactions with other nitrogen compounds make NO dependent upon a

number of species. Because reactions among those species are fast, it is con-

venient to introduce the

odd nitrogen family:

NOx = N + NO + NO2 +

NO 3

+ 2N205 +

HNO4,

(17.11)

the members of which may be thought to remain in photochemical equilibrium

with one another. The lifetime of NOx is much longer than the timescale for

advection (Fig. 17.3), despite the much shorter lifetimes of some of its mem-

bers. The abundance of NO then follows from [NOx] and from partitioning

within the family, which reacts on a slower timescale with other families such

as Ox.

17.2.2 Chlorofluorocarbons

Free radicals of chlorine represent another important link to human activities.

Atomic chlorine is produced naturally in the stratosphere through photodisso-

ciation of methyl chloride CH3C1, which is a product of ocean processes, and

17.2 Involvement of Other Species 543

through its oxidation with reactive species. Atomic chlorine is also produced

by photodissociation of chlorofluorocarbons (CFCs), the sole source of which

is industrial. Like N20, CFC-11 (CFC13) and CFC-12 (CF2C12) are long-lived

(Fig. 17.5). Photochemical lifetimes of several years along with their water in-

solubility, which circumvents normal scavenging mechanisms, allow CFCs to

become homogenized in the troposphere (Fig. 1.20).

In the stratosphere, CFC-11 and CFC-12 decrease upward due to photodis-

sociation, which releases free chlorine in the reactions

CFC13 + hv -+ CFC12 + C1,

(17.12.1)

CF2C12 + hv -+ CF2C1 + C1, (17.12.2)

which operate at wavelengths shorter than 225 nm. Once produced, free CI

reacts with ozone in the catalytic cycle involving chlorine monoxide CIO

C1 + 03 ~ CIO + 02, (17.13.1)

C10 + O --+ CI + 02, (17.13.2)

which has the net effect

0 3 + O --+ 20 2 (net). (17.13.3)

This closed cycle leaves C1 + C10 unchanged, so one atom of chlorine can

destroy many atoms of ozone.

Like free radicals of nitrogen, atomic chlorine can react with other species

on widely varying timescales. It is therefore convenient to introduce the odd

chlorine family:

C1 x = C1 + C10 + HOC1, (17.14)

1 Tin

1 hour 1 day 1month 1year

i I I

i I I I I I I I I I I I I

I I I I I I I I ili I I i I I I I I I i I I ili I [

100 101 102 103 104 105 106 107 108 109

LIFETIME (s)

Figure

17.5 Photochemical lifetimes of chlorine species, including CFC-11

(CFC13)

and

CFC-12 (CF2C12). Source: Brasseur and Solomon (1986).

544

17 The Middle Atmosphere

from which the abundance of C1 follows with the partitioning of member

species. Relative to air motion, CII is short-lived (Fig. 17.5), so it is not pas-

sively advected. Nevertheless, Clx is much longer-lived than its members, which

facilitates the numerical treatment of chlorine species. Odd chlorine can in-

teract with odd nitrogen in the reaction

C10 + NO ~ C1 +

NO2,

(17.15)

which regenerates free C1 and couples the families NOx and Clx. Other inter-

actions among Ox, NOx, and C1 x bear importantly on ozone depletion in the

polar stratosphere (Sec. 17.7).

17.2.3 Methane

Methane enters the equation for ozone because it interacts with NOx and Ox.

In addition to being radiatively active,

CH 4

represents an important link be-

tween chemical constituents in the stratosphere and water vapor, both of which

have the troposphere as a reservoir. Like

N20 ,

methane has a photochemical

lifetime of several years below 30 km (Fig. 17.6) and is water insoluble. These

properties allow

CH 4 to

become well mixed in the troposphere (Fig. 1.20) and

make it useful as a tracer in the stratosphere. Figure 17.7 shows the zonal-

mean distribution of rCH 4 in the middle atmosphere. Mirroring behavior in the

N20

distribution, a dome of high values above the tropical tropopause reflects

upwelling motion in the equatorial stratosphere.

Decreasing mixing ratio in the stratosphere follows from chemical destruc-

tion of

CH 4.

Methane is oxidized by the hydroxyl radical in the reaction

CH 4 --I-OH ~ CH 3 --I- H20.

(17.16.1)

1day 1month 1year lOyrs lOOyrs

80 i I ~ "

g 60

2 4o

I~OH

20 I~CH 4

I ' 104~110 I I [ I I [ ' ' I ' ' I ' '

10 0 101 1 5 10 6 10 7 10 8 10 9 1010

LIFETIME (s)

Figure

17.6 Photochemical lifetimes of hydrogen species. Source: Brasseur and Solomon

(1986).

17.2 Involvement of Other Species

545

..Q

(/)

100

ii!!iiiiiiiiiii!ii!il

!ii!iii!iiiiiiii!!i!

!~i~!i!ii!iiiiiii!~

2.0

1.8

1.6

1.4 E

1.2

1.0

.=_

o.a

0.6

0.4

0.2

0.0

-64 -48 -32

-16

0 16 32 48

Latitude

Figure 17.7 Zonal-mean mixing ratio of methane CH4, as observed by the HALOE instru-

ment on board UARS during September 21-October 15, 1992. Adapted from Bithell

et al.

(1994).

It is also destroyed by atomic oxygen and chlorine in the reactions

CH 4 + O ~ CH 3 + OH,

CH 4 + C1 --+ CH 3 + HC1.

(17.16.2)

(17.16.3)

The latter represents an important sink of reactive Cl x. The free radical CH 3

produced in reactions (17.16) immediately combines with molecular oxygen to

form CH302. Two reactions involving NO~,

CH302 + NO --~ CH30 + NO 2

NO 2 +

hv --+

NO + O,

(17.17.1)

(17.17.2)

then constitute a closed cycle that has the net effect

CH302 +

hv ~

CH30 + O

(net).

(17.17.3)

546

The Middle Atmosphere

10 0

101

Or)

ILl

n'-

12.

102

-90

r. o

(ppmv)

-60 -30 0 30 60 90

LATITUDE

Figure

17.8 Zonal-mean mixing ratio of water vapor observed by Nimbus-7 LIMS during

December, 1978. Adapted from Remsberg

et al.

(1984).

Methane oxidation thus leads to production of Ox. Since the partitioning of Ox

fixes the abundance of 03, reactions (17.17) provide a potentially important

source of ozone in the lower stratosphere.

The free radical OH that initiates the foregoing process is produced by

dissociation of water vapor

H20 + O --+ 2OH.

(17.18)

Above the tropopause, the lifetime of

H20

is several years (Fig. 17.6), so it

too behaves as a tracer at these heights. Water vapor mixing ratio decreases

sharply in the troposphere to a minimum of just a few ppmv at the

hygropause

(Fig. 17.8). The steep decrease with height in the troposphere reflects precip-

itation processes that dehydrate air reaching the tropopause by limiting rH2 o

to saturation values. Above the hygropause, rH~O increases due to oxidation

of methane, as rCH 4 decreases. How much of the observed increase is actually

explained by reactions (17.16) remains unclear.

17.3 Air Motion

Photochemical considerations alone predict maximum column abundances of

ozone at low latitude. This shortcoming is unchanged by the addition of other

chemical species because UV radiation and hence photolysis maximize in the

tropics. Consequently, the maximum of observed ~o3 at middle and high lati-

tudes can be explained only by incorporating dynamical considerations.

17.3 Air

Motion

547

17.3.1 The Brewer-Dobson Circulation

The latitudinal gradient of heating in combination with geostrophic equilib-

rium favors motion that is chiefly zonal, with only a small meridional com-

ponent to transfer heat and chemical species between the equator and poles

(Chapter 15). In the winter hemisphere, ozone heating establishes an equa-

torward temperature gradient over a deep layer (Fig. 1.7), which, by thermal

wind balance, produces strong westerly flow. Stratospheric westerlies intensify

upward along the

polar-night terminator,

where SW absorption vanishes, to

produce the polar-night jet (Fig. 1.8) and the circumpolar vortex (Fig. 1.10b).

In the summer hemisphere, a poleward gradient of heating, which follows from

the distribution of daily insolation (Fig. 1.28), produces a deep temperature

gradient of the opposite sense and strong easterly circumpolar flow.

Under radiative equilibrium, the circulation is zonally symmetric and ex-

periences no net heating. Then, by the first law, individual air parcels do not

cross isentropic surfaces. This implies no net vertical motion and, by continu-

ity, no net meridional motion--analogous to the zonally symmetric circulation

in Sec. 15.2. Therefore, mechanical disturbances that drive the circulation out

of radiative equilibrium play a key role in producing the gradual meridional

overturning that accompanies strong zonal motion in the middle atmosphere.

The behavior of long-lived tracers (Figs. 17.4 and 17.7) implies upwelling

in the tropics, where tropospheric air enters the stratosphere. Downwelling at

middle and high latitudes, which is required by continuity, can then explain the

large abundances of ozone observed in extratropical regions. Air drawn pole-

ward and downward from the chemical source region of ozone in the tropical

stratosphere (Fig. 17.9) would then undergo compression to yield the greatest

absolute concentrations

/903

(Fig. 1.17) and hence the greatest column abun-

LI.I

I--

116 I

106

I ! I I I I

90S 60 50 0 50 60 90N

LATITUDE

Figure

17.9 Streamlines of the mean meridional circulation of the middle atmosphere, in

a quasi-Lagrangian representation. Adapted from Garcia and Solomon (1983), copyright by the

American Geophysical Union.

548

17 The Middle Atmosphere

dances at middle and high latitudes. Suggested by Dobson (1930) and Brewer

(1949) to explain tracer observations, this gradual meridional overturning is

responsible for observed distributions of ozone and other chemical species in

the middle atmoSphere. The Brewer-Dobson circulation also underlies chemi-

cal production because, by displacing air out of photochemical equilibrium, it

enables chemical reactions to occur.

17.3.2 Wave Driving of the Mean Meridional Circulation

Although it may be envisioned as a zonally symmetric overturning in the

meridional plane, the Brewer-Dobson circulation is in reality accomplished

by zonally asymmetric processes. A clue to its origin lies in the radiative-

equilibrium state of the middle atmosphere (Fig. 17.10). During solstice, the

sharp gradient of heating across the polar-night terminator produces radiative-

equilibrium temperatures TRE colder than 150 K (Fig. 17.10a). These are much

colder than observed zonal-mean temperatures (Fig. 1.7), which are more

uniform latitudinally and remain above 200 K in the Northern Hemisphere.

Radiative-equilibrium winds implied by the deep layer of sharp temperature

gradient (Fig. 17.10b) intensify to more than 300 m s-l--much stronger than

observed winds (Fig. 1.8).

The observed polar-night vortex is warmer than radiative equilibrium, so it

must experience net radiative cooling (Sec. 8.5). Observed temperatures and

ozone imply cooling rates exceeding 8 K day -1 in the polar night (Fig. 8.27). By

the first law, that cooling must reduce the potential temperature of air parcels

inside the vortex. Positive stability (30/3z > 0) then implies downwelling across

isentropic surfaces (e.g., parcels sink to lower 0). Calculations in which this

diabatic cooling is prescribed (Murgatroyd and Singleton, 1961) qualitatively

reproduce the meridional circulation inferred from tracer behavior by Brewer

and Dobson.

In the Southern Hemisphere, polar temperatures as cold as 180 K are

observed. Zonal-mean winds are then commensurately faster. Colder and

stronger than its counterpart over the Arctic, the Antarctic polar-night vortex

remains closer to radiative-equilibrium behavior. According to the preceding

arguments, radiative cooling and downwelling in extratropical regions should

then be weaker. Similar considerations apply to the summertime circulation

(compare Fig. 8.27), in which strong easterlies block planetary wave propaga-

tion from below (Sec. 14.5). 2

The key to understanding interhemispheric differences of the circumpolar

vortex lies in understanding how the circulation is maintained out of radiative

equilibrium. Observed motion during winter is rarely in a quiescent state of

zonal symmetry, especially in the Northern Hemisphere, where the circumpolar

2The absence of eddy motion enables the cloud of volcanic debris in Fig. 9.6 to remain

confined meridionally through the summer season. When the zonal flow reverses near equinox and

planetary waves emerge from below, the volcanic cloud is quickly dispersed over the hemisphere.

17.3 Air

Motion

549

iaJ

I---

(a)

-90 ~ -60 o -30 o 0 o 30 ~ 60 ~

LATITUDE

(b)

llll l//i

I I I IllU I!

', ', ',',',:~ A

t t t ii!11 ~

I I I IIII ~mll!

I I u II11 IiIg I

I I I~..11

~- ' , ~u'J'.elilNI

! % 200 V llll~

I 1 I III " i iP

g 80 /

,, ,, ,,,,,,,,/~ilillll

I A \ \\~'////

..... .~:.~lll%//\ \ \ k.;///~

I I I I I II I!

I I I I| II I|

', ',',

'' ''~

o',',',

/ t

,, ,, ,., !,,0\\ "---J /

20 I l "

90*

-90 ~ -60 o -30 o 0 o 30 ~ 60 ~ 90 ~

LATITUDE

Figure 17.10 Radiative-equilibrium (a) temperature and (b) zonal wind in the middle at-

mosphere during solstice, as calculated in a radiative-convective-photochemical model. Thermal

structure adapted from Fels (1985). Zonal wind calculated from thermal wind balance and from

climatological motion at 20 km in Fig. 1.8.

flow is continually disturbed by planetary waves that propagate upward from

the troposphere. Planetary waves in the Northern Hemisphere are stronger

than those in the Southern Hemisphere because of larger orographic forcing

at the earth's surface. Able to penetrate into the strong westerlies of the winter

stratosphere, planetary waves drive the circulation out of zonal symmetry and

away from radiative equilibrium. As illustrated by Fig. 1.10a, air then flows